1. Introduction
1.1 Motivation and Objectives
Ice storms are among the most hazardous, disruptive, and costly winter weather
phenomena. From a socioeconomic perspective, ice storms endanger human life and
safety, undermine public infrastructure, and adversely impact local and regional
economies. During the 5–9 January 1998 North American ice storm, prolonged freezing
rain resulted in widespread power outages, human fatalities and injuries, and catastrophic
damage to forests and agricultural operations throughout northern New York, New
England, and southeastern Canada (DeGaetano 2000). Power disruption affected nearly
4.5 million people, and the region’s dairy, maple syrup, and timber industries all suffered
considerable losses. Insurance claims totaled at least $175 million across New York,
Vermont, New Hampshire, and Maine. Despite common public opinion that the 1998 ice
storm was an unprecedented event, New York and New England have endured similar
magnitude events since the early twentieth century (DeGaetano 2000; Gyakum and
Roebber 2001).
More recently (2008), a severe ice storm in south-central China caused $22.3
billion in direct economic losses, 129 human fatalities, and power failure and structural
damage that displaced 1.7 million people (Zhou et al. 2011). Electrical power disruption
and heavy ice accretion compromised the nation’s infrastructure and jeopardized the
distribution of food and other necessities. Extensive forest damage eventually led to soil
erosion, landslides, insect infestation, and numerous forest fires. The Chinese ice storm
exemplifies the socioeconomic and ecological consequences that can result from
unsustainable practices such as concentrated industrial development and short-term
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growth policies (Zhou et al. 2011). Call (2010) argues that a growing dependence on
electricity has also increased the magnitude and duration of ice storm-related power
outages, thereby elevating societal vulnerability to costly disasters.
From a geographical perspective, ice storms are historically prevalent and
destructive across the United States, particularly in the Northeast. Changnon (2003a)
identified 87 ice-storm catastrophes during the 1949–2000 period that caused an
estimated $16.3 billion in insured losses throughout the contiguous U.S. As illustrated by
Figs. 1.1 and 1.2, the Northeast region experienced the highest frequency of ice storm
catastrophes (39) and suffered the greatest insured losses (roughly $4.84 billion total;
$124 million per event). More than 20 ice-storm catastrophes occurred in nine states,
including North Carolina, Virginia, Maryland, Pennsylvania, New Jersey, New York,
Connecticut, Rhode Island, and Massachusetts. Changnon (2003a) notes that locations
along and east of the Appalachian Mountains are climatologically favorable for cold-air
damming and low-level frontal development, both of which are commonly observed in
association with freezing rain.
A separate 9-cool-season (1928–1937) database of ice accretion measurements
taken by railroad personnel (Hay 1957) also revealed a maximum in the frequency of
damaging ice-storm areas in the northeastern U.S. National Weather Service (NWS)
Warning Coordination Meteorologists (WCMs) estimate that ice storm recurrence
intervals range between 0.5 and 2 years throughout the NWS Eastern Region, with the
shortest recurrence intervals observed across northern and interior sections (Call 2009).
The geographical distributions of ice storms from Changnon (2003a) are consistent with
studies on the climatological frequency of freezing rain. For instance, Houston and
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Changnon (2007) determined that the greatest annual average number of freezing rain
hours occurs in upstate New York and New England, with an additional distinct
maximum east of the southern Appalachians (Fig. 1.3).
From a meteorological perspective, ice storms present a major operational
forecast challenge due to the combined influence of synoptic, mesoscale, and
microphysical processes on precipitation type. Based on a 2006 survey completed by 15
WCMs in the NWS Eastern Region, most offices would issue watches 24–48 h preceding
an ice storm, but only eight (53%) would issue warnings more than 24 h in advance (Call
2009). These responses convey how the difficulty in forecasting precipitation type and
intensity undermines the ability to issue precise forecasts beyond 24 h. Moreover,
empirical evidence suggests that operational models and forecasters often underestimate
the extent and duration of freezing rain in complex winter storms.
In recent years, researchers have developed precipitation type algorithms utilizing
parameters such as wet bulb temperature, surface temperature, relative humidity, ice
fraction, and cold/warm layer depth (Wandishin et al. 2005). Despite our knowledge of
the meteorological conditions and physical mechanisms that influence precipitation type,
we cannot accurately observe nor model critical thermodynamic variables and
microphysical processes at sufficiently high spatial and temporal resolutions to precisely
forecast precipitation type. Wandishin et al. (2005) nevertheless showed that ensemble
forecasts incorporating various precipitation-type algorithms can increase short-range (0–
48 h) forecast skill. Utility companies, emergency managers, and the general public
would all benefit from improved short-to-medium range prediction of freezing rain (Call
2010). As DeGaetano et al. (2008) note, utility companies encounter many challenging
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decisions before an impending ice storm, and the lack of forecast tools for predicting ice
accretion further complicates these decisions. Thus, adequate preparation and risk
management require that accurate forecasts of icing amount and location be issued with
ample lead time.
In consideration of socioeconomic and forecast issues highlighted above, we
outline three primary objectives for this study. First, we seek to create long-term
climatologies of freezing rain and ice storms in the northeastern U.S. These climatologies
will examine the temporal and spatial variability of freezing rain and ice storms, as well
as characterize ice storms based on spatial properties and relevant meteorological
features. Second, we aim to identify the antecedent environments conducive to ice storms
and the physical mechanisms that govern their evolution. Although freezing rain typically
occurs under preferred synoptic conditions, we need to consider how mesoscale processes
modify dynamically and thermodynamically forced large-scale circulations and
associated quasi-geostrophic (QG) forcing on regional and local scales. Third, we hope to
increase situational awareness and provide scientific insights that will ultimately improve
the future prediction of ice storms. The anticipated findings will build upon existing
conceptual models and help facilitate critical decisions made by operational forecasters,
utility companies, and emergency managers.
1.2 Background
1.2.1 Microphysical Processes and Precipitation Type Issues
Although this study focuses on the synoptic and mesoscale dynamics associated
with freezing rain, it is worth noting how microphysical processes and the vertical
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distribution of thermodynamic variables ultimately determine precipitation type and
evolution. Bocchieri (1980) analyzed rawindsonde observations from 48 upper-air
stations to objectively identify which physical parameters most influenced precipitation
type. Figure 1.4 shows composite temperature and dewpoint profiles for 94 freezing rain
and 127 freezing drizzle observations. The classical freezing rain sounding is
characterized by a nearly saturated vertical profile, a surface-based subfreezing layer, and
a distinct melting layer (T > 0°C) aloft, whereas freezing drizzle typically occurs in
colder, drier environments lacking well-defined warm layers. Based on his statistical
analysis, Bocchieri proposed using six critical parameters to diagnose precipitation type:
1) the mean layer temperature in the lowest 1000 m, 2) the mean layer temperature
between 500 m and 2500 m, 3) the warm layer depth, 4) the warm layer area (bounded by
the temperature profile and the 0°C isotherm), 5) the cold layer depth, and 6) the cold
layer area (bounded by the wet-bulb temperature profile and the 0°C isotherm). Although
this method effectively discriminated between liquid and frozen precipitation, it failed to
accurately predict freezing precipitation. Errors likely arose from the fact that Bocchieri
neglected the impacts of microphysical processes on precipitation type.
Stewart and King (1987) examined the mesoscale substructure and precipitation
evolution in three midlatitude winter storms. Freezing rain occurred in a variety of
environments, including warm frontal bands, bands along the cold front, and bands
trailing the low center. Model simulations demonstrated that initial snowflake size,
precipitation rate, and the depths of the melting and refreezing layers largely determined
whether precipitation fell as freezing rain/drizzle or ice pellets. In addition, relative
humidity can influence hydrometeor phase changes via diabatic effects (evaporational
5
cooling). Their results suggested that smaller particles, lower precipitation rates,
extensive (shallow) melting (refreezing) layers, and saturated conditions increase the
likelihood of freezing rain versus ice pellets.
Raga et al. (1991) investigated the relationship between microphysical processes
and thermodynamic and kinematic properties of the transition region in a midlatitude
winter storm. Aircraft observations indicated an enhanced horizontal temperature
gradient at the cold edge of the transition region, likely driven by diabatic warming
associated with the freezing of cloud droplets and partially melted particles below the
melting layer. The enhanced meridional temperature gradient coincided with an increase
in vertical wind shear and the formation of a low-level easterly jet. Observations also
revealed strong southerly flow rising above the subfreezing air at low levels and
transporting warm, moist air into the transition region. These easterly and southerly jets
correspond to the cold and warm conveyor belts, respectively, noted by Browning (1986).
As illustrated by Fig. 1.5, Raga et al. (1991) developed a conceptual diagram highlighting
the predominant microphysical processes, thermodynamic features, and kinematic
perturbations associated with transition regions.
Szeto and Stewart (1997) evaluated the effects of melting on surface
frontogenesis and related thermodynamic and kinematic fields. Transition regions are
often located near low-level frontal boundaries, and microphysical processes can modify
these boundaries via latent heating/cooling. Model simulations suggested that cooling by
melting accelerates low-level frontogenesis, induces downdrafts and outflows that
enhance convergence and ascent on both sides of the transition region, and increases
vertical wind shear in the along-front and cross-front directions. These results are
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consistent with Szeto et al. (1988), who concluded that melting-induced cooling triggers a
thermally indirect mesoscale circulation centered near the cold edge of the rain–snow
boundary. Furthermore, in the presence of large vertical wind shear, cooling by melting
will oppose differential temperature advection and thereby influence the evolution of the
vertical temperature profile.
Zerr (1997) examined 34 soundings and described the microphysical and
thermodynamic features associated with freezing rain versus ice pellets. In general,
shallow melting layers, deep refreezing layers, colder profiles, and subsaturated
conditions appeared to inhibit complete melting and thus prevent freezing rain from
occurring. Zerr also defined a melting parameter (βM = TmaxΔZM) and a refreezing
parameter (βF = TminΔZF), where Tmax (Tmin) is the maximum (minimum) temperature in
the melting (refreezing) layer, and ΔZM (ΔZF) is the depth of the melting (refreezing)
layer. Refreezing parameters were large regardless of precipitation type, but melting
parameters were notably smaller during ice pellet events than during freezing rain events.
Most soundings were characterized by veering winds and warm advection aloft, but
several soundings indicated backing winds and low-level cold advection (likely
accompanied by a cold frontal passage).
Rauber et al. (2000) compared the relative importance of warm cloud rain and
melting processes during freezing precipitation events. Freezing precipitation can occur
via the classical “melting process” (ice particles undergo melting in an above-freezing
layer and become supercooled after passing through a surface-based subfreezing layer) or
the “warm rain process” (cloud droplets undergo collision and coalescence in subfreezing
air that lacks a melting layer). The authors determined that the warm rain process was
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likely active in about 75% of the 972 analyzed soundings, but typically resulted in
freezing drizzle. On the contrary, only 25% of soundings were characterized by
conditions favorable for the classical melting process (cloud top temperatures < −10°C
and a warm layer aloft). Soundings of the latter type were most frequently observed
during freezing rain events in the Midwest U.S. and east of the Appalachians.
Rauber et al. (2001a) assessed the effectiveness of a nondimensional parameter
(Czys et al. 1996) that distinguishes between freezing rain and ice pellets. Defined as τ =
tres
tmelt
, this parameter uses values of warm layer depth, ambient temperature, particle fall
speed, and critical ice particle radius to estimate the ratio of residence time in the warm
layer (tres) to the time required for complete melting (tmelt). Employing a 25-year
climatology of soundings during freezing rain, freezing drizzle, and ice pellet events,
Rauber et al. (2001a) exposed two key issues with this parameter. First, only 306 of the
1052 soundings conformed to the vertical profile (cold cloud tops, elevated warm layers,
and subfreezing surface layers) considered by Czys et al. (1996). Second, this parameter
poorly discriminated between freezing rain/drizzle and ice pellets.
Theriault et al. (2010) employed a one-dimensional kinematic cloud model to
simulate the behavior of hydrometeors and evaluate precipitation type sensitivity to both
temperature and precipitation intensity. Differentiating between freezing rain and ice
pellets is especially difficult because supercooled droplets entering the subfreezing layer
may remain unfrozen, undergo refreezing, or interact with locally generated ice crystals.
Their sensitivity analysis demonstrated that higher precipitation rates and larger initial
snowflakes would inhibit completing melting. Slight changes (±0.5°C) in the temperature
profile would also have major implications for precipitation type (warmer profiles favor
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complete melting and, consequently, freezing rain). Since precipitation type at the surface
strongly depends on temperature, precipitation intensity, and poorly resolved
microphysical processes, precise forecasts of precipitation type are a likely major
weakness of operational models.
1.2.2 Climatological Perspective
Numerous studies have examined the climatological aspects of freezing rain and
ice storms across North America. Branick (1997) constructed a 13-year (1982–1994)
climatology of significant winter weather events impacting the contiguous U.S. He
determined that 25% of all events featured icing, but only 12% met the 0.25 in (0.64 cm)
minimum ice storm criterion. Approximately, 84% of all events impacted areas less than
250,000 km2, and 70% of precipitation events lasted between 6 and 24 h locally. These
spatial and temporal characteristics suggest that although ice storms typically occur in
preferred synoptic settings, they are predominantly mesoscale phenomena.
Bernstein (2000) investigated local and regional influences on the type of freezing
precipitation (freezing rain, freezing drizzle, and ice pellets) at various sounding sites in
the continental U.S. The 30-year (1961–1990) climatology revealed high freezing
precipitation frequencies (> 30 h of freezing precipitation per year) along and east of the
Appalachian Mountains, where freezing rain represented 30–50% of all freezing
precipitation. In the lee of the southern Appalachians, north-northeasterly surface winds
and a meridional pressure dipole (higher pressure to the north and lower pressure to the
south) are consistent with cold-air damming. Across New England, wind directions
varying between northeasterly and northwesterly suggest the importance of both cold-air
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damming and frontal boundaries enhanced by land–ocean thermal contrasts. Cold cloudtop temperatures (< −10°C), well-defined melting layers, and relatively shallow
subfreezing layers in both regions generally support freezing rain formation via the
melting process.
Focusing on the 1976–1990 period, Cortinas (2000) evaluated the spatial and
temporal distribution of freezing rain throughout the Great Lakes Region, and identified
synoptic-scale features characteristic of freezing rain events. Freezing rain occurred more
frequently across the eastern Great Lakes than across the western Great Lakes, with
regional maxima over interior sections of southeastern Ontario, west-central
Pennsylvania, and upstate New York, and regional minima along the western lake shores.
These spatial variations are likely influenced by: 1) the frequency of surface cyclone
tracks and associated frontal boundaries, 2) availability of Atlantic moisture, and 3) local
and regional topography. Moreover, the notable decrease in frequency near the lake
shores implies that large bodies of water can modify the thermodynamic environment and
ultimately reduce the likelihood of freezing rain. Freezing rain reports exhibited
significant diurnal variability, with the highest (lowest) frequencies observed during the
morning (afternoon). Most events were short-lived and accompanied by a transient
surface cyclone tracking east-northeast through the Midwest, anomalously high pressure
over eastern North America, and easterly surface winds.
Rauber et al. (2001b) discussed common synoptic-scale features associated with
411 freezing precipitation events east of the Rocky Mountains during the 1970–1994
period. Based on the locations and orientations of surface cyclones, anticyclones, and
frontal boundaries, the authors defined seven archetypical patterns conducive to freezing
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precipitation. Freezing precipitation occurred most frequently on the poleward side of a
warm front or within the occluded region of a surface cyclone. Cold-air damming (east of
the southern and central Appalachians) and cold-air trapping (along the northern
Appalachians) were also present in a substantial number of freezing precipitation events.
Surface winds varied between easterly and northwesterly, whereas winds aloft were
predominantly southwesterly.
Adopting the same period (1976–1990) as Cortinas (2000), Robbins and Cortinas
(2002) analyzed the local and synoptic environments associated with freezing rain events
in different regions. Approximately 68% of freezing rain events in the Piedmont region
(interior North Carolina and Virginia) were associated with cold-air damming east of the
Appalachians. Low-level northeasterly winds and cold advection established a thermal
trough and pressure ridge at the surface, while moist, southerly flow and warm advection
produced a warm layer aloft. Most freezing rain events in the Allegheny–Catskill region
(west-central Pennsylvania and east-central New York) were associated with a surface
cyclone passing to the west, minimal thermal advection at the surface, and pronounced
warm advection aloft via southwesterly flow. The composite surface wind field also
reveals a second area of cyclonic vorticity near the mid-Atlantic coast, and thereby
suggests that some events may occur in conjunction with Miller Type A (cyclogenesis
near the Gulf Coast) or Miller Type B (secondary cyclogenesis along the East Coast)
cyclone tracks (Miller 1946). Roughly 57% of all freezing rain events occurred along a
stationary or quasi-stationary warm front, and 33% occurred within the northern,
northwestern, or western sector of surface cyclones.
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Employing a 9-cool-season (1928–1937) record of glaze accumulations from the
Association of American Railroads (Hay 1957), Changnon (2003a) evaluated the
geographical distribution, spatial properties, and ice thicknesses associated with 368
damaging ice-storm areas. In the northeastern U.S., these ice-storm areas were
characterized by elongated shapes (length-to-width ratios > 2:1) and either southwest–tonortheast or south-southwest-to-north-northeast orientations. Radial ice thicknesses
averaged 1.0 cm (0.39 in) across the region, with 25% of measurements exceeding 2.0 cm
(0.79 in). Consistent with Branick’s (1997) significant winter weather climatology, the
number of ice storms and spatial coverage were inversely related.
Utilizing a long-term (1945–2000), high-resolution database, Changnon (2003b)
assessed the urban modification of freezing rain near four selected U.S. cities. Compared
to surrounding stations, urban stations within New York, NY, and Chicago, IL,
experienced 16–43% fewer freezing rain days, while urban stations within Washington,
DC, and St. Louis, MO, experienced 9–30% fewer freezing rain days. In New York and
Chicago, both urban heat-island effects and maritime influences (lake/ocean proximity)
were responsible for the observed reduction in freezing rain frequency.
Changnon and Karl (2003) examined the spatial and temporal variability of
freezing rain days in the contiguous U.S. during the 1948–2000 period. On average, more
than 4 freezing rain days occur annually in a west-southwest-to-east-northeast band
extending from Missouri to central New York and a south-southwest-to-north-northeast
band between the Appalachian Mountains and the Atlantic coast. Freezing rain events in
the Midwest and interior Northeast primarily involve warm advection aloft and
overrunning precipitation on the cold side of surface frontal boundaries. Locations along
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and east of the Appalachians often experience freezing rain in association with a warm
maritime air mass displaced above a shallow region of cold-air damming. Throughout the
eastern U.S., the monthly frequency of freezing rain days is greatest during January,
while the yearly frequency exhibits significant interannual and interdecadal variability.
Cortinas et al. (2004) analyzed the spatial distribution, temporal variability, and
surface conditions associated with freezing rain, freezing drizzle, and ice pellets in the
U.S. and Canada. During the 1976–1990 period, freezing rain was most prevalent across
eastern North America, particularly over the interior northeastern U.S., the Canadian
Maritimes, southeastern Quebec, and Newfoundland. The observed geographical patterns
likely result from several factors, including topography, proximity to large bodies of
water, and midlatitude cyclone activity. In the contiguous U.S., freezing rain and freezing
drizzle predominately occur between December and March. Freezing rain was generally
short-lived, strongly tied to the diurnal heating cycle, and coincident with surface
temperatures near or slightly below 0°C.
1.2.3 Synoptic and Mesoscale Dynamics
Besides recognizing which geographical regions and meteorological conditions
are climatologically favorable for ice storms, we need to understand the dynamical
mechanisms that modulate the occurrence and duration of freezing rain. Forbes et al.
(1987) investigated synoptic-scale and mesoscale circulations and thermal patterns
associated with cold-air damming during an Appalachian ice storm. Before the icing
occurred, an anomalously strong surface anticyclone approached the northeastern U.S.,
and the resulting pressure gradient induced easterly and northeasterly geostrophic winds
13
along and east of the Appalachians. Meanwhile, high static stability and frictional effects
restricted airflow over the mountains and triggered northerly ageostrophic winds and a
mountain-parallel jet at roughly 500 m. Low-level ageostrophic cold advection, in
concert with geostrophic warm advection downstream of an 850-hPa short-wave trough,
established a distinct surface-based cold dome and pressure ridge below a deep inversion
layer. Consequently, warm, moist air from the Atlantic Ocean was displaced above the
cold dome and resulted in overrunning precipitation along the eastern slopes of the
southern Appalachians.
Rauber et al. (1994) analyzed the synoptic and mesoscale structure of the poorly
forecasted St. Valentine’s Day ice storm in north-central Illinois. During this event, a
narrow swath of heavy freezing rain occurred on the poleward side of a zonally
elongated, quasi-stationary warm front. South-southwesterly flow extending from the
Gulf of Mexico to the Midwest provided strong low- to midlevel warm advection and
moisture transport, and produced a saturated inversion layer above a shallow but
persistent subfreezing layer. The highest freezing rain amounts fell as warm, moist air
was lifted above the warm frontal boundary. In addition, the authors argued that melting
and sublimation of accumulated ice reinforced the meridional temperature gradient,
thereby inhibiting the northward progression of the surface warm front and ultimately
prolonging the freezing rain event.
Szeto et al. (1999) utilized a cloud-resolving model to diagnose the mesoscale
processes governing the evolution of an ice storm across eastern Canada. As the
deepening cyclone and associated warm front approached the Canadian Maritimes and
Newfoundland, sharp land–ocean contrasts in surface friction and temperature led to the
14
intensification of the ageostrophic cross-frontal circulation. In turn, the enhanced crossfrontal circulation accelerated frontogenesis and strengthened the vertical differential
temperature advection, thereby creating a low-level inversion on the poleward side of the
surface warm front. Furthermore, increased surface convergence along the warm front
contributed to sloped ascent and overrunning precipitation. Frozen precipitation
originating within bands above the sloping inversion melted as it fell through the warm
layer and, depending on the depth of the subfreezing layer, reached the surface either as
ice pellets or freezing rain.
Gyakum and Roebber (2001) described how the large-scale circulation and
thermodynamic environment fostered excessive freezing rain during the 5–9 Jan 1998 ice
storm. Composites of five-day mean sea level pressure, 1000−500-hPa thickness, and
1000−850-hPa thickness revealed an inverted trough extending from the Gulf Coast to
the eastern Great Lakes, a strong anticyclone near James Bay (Figs. 1.6a and 1.6b), an
anomalous thickness ridge over the eastern U.S., and an enhanced baroclinic zone across
southeastern Canada (Figs. 1.6c and 1.6d). Despite remarkable warm anomalies
throughout much of the troposphere, persistent northeasterly surface winds maintained a
shallow layer of subfreezing air in the ice storm region. Using an Eulerian moisture
budget analysis, the authors determined that moisture convergence and advection were
largely responsible for enhancing the magnitude and longevity of precipitation. This
event was unique in that most locations received freezing rain in multiple distinct periods
as a series of short-wave disturbances tracked through the Ohio Valley and eastern Great
Lakes. As indicated by the backward trajectories in Fig. 1.7, air parcels reaching the
precipitation zone during the first (second) episode underwent substantial moistening and
15
latent heating over the Gulf of Mexico (Atlantic Ocean). Heavier precipitation in the
second period resulted from air parcels having much longer exposure to the subtropical
Atlantic and thus attaining higher mixing ratios (12.2 g kg−1 versus 10.5 g kg−1) and
equivalent potential temperatures (330 K versus 320 K).
Roebber and Gyakum (2003) assessed the orographic modification of low-level
winds, frontogenesis, and precipitation amounts during the 5–9 Jan 1998 ice storm.
Throughout the event, the presence of an anomalously strong surface anticyclone over
central Quebec and lower pressure to the southeast induced orographic channeling within
the St. Lawrence and Champlain Valleys. In the southwest-to-northeast oriented St.
Lawrence Valley, pressure-driven channeling resulted in persistent northeasterly winds
that continually reinforced cold air near the surface. On the contrary, low-level winds in
the north–south oriented Champlain Valley were quite sensitive to changes in the
horizontal pressure gradient and alternated between south-southwesterly and northnortheasterly. Ageostrophic cold advection in the St. Lawrence Valley and geostrophic
warm advection to the south also provided a frontogenetical focus that regionally
enhanced precipitation. Based on numerical model simulations, the authors concluded
that freezing rain amounts would have been significantly lower in the absence of these
topographic features.
Ramos da Silva et al. (2006) evaluated the sensitivity of freezing rain to Atlantic
sea surface temperatures (SSTs). Citing the 4–5 Dec 2002 case as an example, the authors
noted that ice storms east of the Appalachians typically occur within a region of cold-air
damming, beneath a deep inversion produced by moist onshore flow and warm advection
aloft. Numerical model simulations of 27 ice storms during the 1991–2004 period
16
revealed a significant positive correlation between melting layer depth and Atlantic SSTs.
An increase in Atlantic SSTs would likely enhance the advection of warm, moist air
above the damming region, create a deeper, stronger, and more persistent melting layer,
and ultimately yield greater freezing rain amounts. Given these results, one may speculate
that Atlantic SSTs also modulate ice storms across New England, where cold-air
damming and coastal frontogenesis (e.g., Bosart 1975) are common during the coolseason.
Sun and Zhao (2010) and Zhou et al. (2011) identified critical synoptic–mesoscale
linkages associated with multiple freezing rain episodes during the 10 Jan–6 Feb 2008
Chinese ice storm. The remarkable longevity of this event was attributed to persistent
large-scale circulation anomalies in the subtropics and midlatitudes. A blocking ridge
over western Siberia provided a continuous source of continental polar air (aided by
regional topographic features), while southwesterly flow and frequent short-wave
disturbances originating within an unusually active southern stream transported maritime
tropical air across southern China (Fig. 1.8). Moreover, the juxtaposition of these two air
masses established a quasi-stationary frontal boundary characterized by an elevated
inversion and saturated conditions below 700 hPa. Moisture convergence, frontogenesis,
and sloped ascent contributed to overrunning precipitation in the form of freezing rain.
Ressler et al. (2012) considered 46 long-duration freezing rain episodes at
Montreal, QC (YUL), and discussed how event duration, precipitation rate, moisture
transport, thermal advection, and QG forcing varied depending on the position and
orientation of the upstream 500-hPa trough axis. All events demonstrated the importance
of low- to midlevel warm advection, poleward moisture transport, and surface cold
17
advection via pressure-driven channeling. As previously noted, these processes were
instrumental in prolonging freezing rain during the 5–9 Jan 1998 ice storm (Gyakum and
Roebber 2001; Roebber and Gyakum 2003). Positively tilted, long-wave troughs over
western North America (west events) were generally associated with longer-duration
events but lower precipitation rates and weaker dynamical forcing for ascent. Negatively
tilted, short-wave troughs over eastern North America (east events) were often associated
with progressive cyclones and shorter-duration events but higher precipitation rates and
stronger dynamical forcing for ascent. During west events, such as the 5–9 Jan 1998 ice
storm, the quasi-stationary nature of the synoptic-scale circulation (higher pressure to the
northeast, lower pressure to the southwest) helped sustain a thermodynamic profile
conductive to freezing rain.
1.3 Thesis Outline
The following section (Chapter 2) will provide an overview of the data and
methodology used in this study. Chapter 3 will examine the climatological aspects of
freezing rain and ice storms across the northeastern U.S. Chapter 4 will utilize a
composite analysis to discuss the large-scale circulation patterns, thermodynamic
environments, QG forcing, and synoptic–mesoscale linkages commonly associated with
ice storms. Chapter 5 will employ a case study analysis to describe the synoptic evolution
of two recent ice storms, as well as the dynamical mechanisms that contributed to
prolonged freezing rain and significant ice accretions. Chapter 6 will expand upon the
results of the climatologies, composite analysis, and case studies, introduce conceptual
models, and conclude with a brief summary of recommendations for future work.
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Figure 1.1: The number of ice-storm catastrophes in each climate region during 1949–
2000. Values in parentheses are those catastrophes that only occurred within the region.
Caption and figure reproduced from Fig. 3 in Changnon (2003a).
Figure 1.2: The amount of loss (millions of dollars expressed in 2000 values) from icestorm catastrophes in each climate region during 1949–2000. Values in parentheses are
the average losses per catastrophe. Caption and figure reproduced from Fig. 2 in
Changnon (2003a).
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Figure 1.3: Average number of hours per year with freezing rain in the United States.
Caption and figure reproduced from Fig. 1 in Houston and Changnon (2007).
Figure 1.4: Composite temperature and dewpoint profiles for freezing drizzle (ZL) and
freezing rain (ZR). The sample consists of 94 freezing rain and 127 freezing drizzle
rawinsonde observations from 48 stations between Oct 1972 and April 1976. Figure
reproduced from Fig. 1 in Bocchieri (1980).
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Figure 1.5: Conceptual model of the transition region in a midlatitude winter storm.
Figure reproduced from Fig. 14 in Raga et al. (1991).
21
Figure 1.6: (a) Time-mean sea level pressure (heavy solid, interval of 4 hPa) and 1000–
500-hPa thickness (dashed, interval of 60 m) for the 5–9 Jan 1998 period. Time-mean
anomalies of (b) sea level pressure [heavy contours with interval of 4 hPa and solid
(dashed) for positive (negative) values], (c) 1000–500-hPa thickness [heavy contours
with interval of 60 m with solid (dashed) for positive (negative) values], and (d) 1000–
925 hPa thickness [heavy contours with interval of 7 m with solid (dashed) for positive
(negative) values]. Thickness anomaly contour intervals in (c) and (d) correspond to
approximately 3°C. Caption and figure reproduced from Fig. 2 in Gyakum and Roebber
(2001).
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Figure 1.7: Paths of three parcel trajectories ending in the precipitation zone (194,400
km2 boxed region) during the 1998 ice storm. Tick marks indicate 12-hourly positions,
beginning with the noted start time. Figure reproduced from Fig. 8 in Gyakum and
Roebber (2001).
Figure 1.8: The encounter of continental polar (cP; blue arrows) with tropical
maritime (mT; red arrows) air masses to the east of Tibetan Plateau. The pink
oval marks the ice storm region. The thicker arrows indicate the dominant
directions in airmass movement. Caption and figure reproduced from Fig. 1 in Zhou et al.
(2011).
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